Throughout the Earth's history, climate (defined
as the long-term statistical behavior of the atmosphere) has been
characterized by numerous cycles with successive fluctuations between colder
and warmer periods. Even though the mean temperature of the Earth over the
past geological periods has probably not varied by more than a few degrees
Celsius, climate changes have produced dramatic variations: in the level of
the oceans, in the geographical extent of the ice sheets, in water supply, and
in the distribution of continental ecosystems, for example. Over the past
centuries, perturbations associated with economic development driven by
growing populations, and specifically with agricultural practices and
industrial activities, have altered our chemical and physical environment with
potential effects on the climate system.

An example is provided by the release to the atmosphere of increasing
quantities of carbon dioxide (CO2) as a result of fossil fuel
consumption and biomass burning. As noticed many years ago (see
Box 15.1), increases in the concentration of
this gas tend to enhance the absorption by the atmosphere of outgoing
terrestrial radiation from the surface, and at the same time to enhance the
infrared radiation emitted at higher altitudes and colder temperatures. The
net effect, often called the "greenhouse effect" is a positive radiative
forcing, which tends to warm the lower atmosphere and the surface. Other
radiatively active gases, such as methane (CH4), nitrous oxide
(N2O), the chlorofluorocarbons (CFCs), and ozone (O3),
the atmospheric abundance of which has changed as a result of human
activities, trap additional radiative energy in the Earth-atmosphere system.

A second example is provided by anthropogenic aerosols present in the
troposphere. These small airborne particles reflect and absorb solar
radiation. Through their effects on cloud properties, they can also alter
cloud microphysical properties, including cloud reflectivity. In most cases
the net effect is a negative forcing, which tends to cool the Earth's climate.

The response of the Earth system to these changes in the radiative balance
is difficult to estimate. Atmospheric and oceanic temperatures as well as
associated weather patterns are expected to be altered. Changes in the
hydrological cycle, and specifically modifications in the precipitation and
evaporation regimes (including drought and flood frequencies), and changes
in cloudiness are also expected to result from modified radiative forcing.

Human-induced effects on climate are and will be superimposed on natural
climate variability. It is therefore difficult to establish with a sufficient
degree of confidence that the observed trend in the globally averaged
temperature of the Earth (0.7 K since the year 1850) is the result of
anthropogenic activity. However, on the basis of the geographical, seasonal,
and vertical patterns in the observed temperature changes, there is growing
evidence of a human influence on the recent evolution of global and regional
climate.

This chapter focuses on the interactions between atmospheric chemistry and
climate and presents the different processes that contribute to the Earth's
radiative balance. The relationship of past climate evolution to variations
in the chemical composition of the atmosphere is also discussed. Finally, we
examine the potential impact of anthropogenic trace gases and aerosols on the
climate system.

15.2 Radiation in the Atmosphere

15.2.1 Solar Radiation

Solar radiation is the primary source of energy for the Earth system. This
energy is provided mostly in the form of ultraviolet, visible, and
near-infrared radiation (wavelength between 0.2 and 4 µm). At the top of
the atmosphere, the shortwave energy flux intercepted by a surface normal to
the direction of the Sun is approximately equal to 1370 W m-2, and
is called the solar constant. The corresponding energy captured by the
Earth's system is on the average 342 W m-2. This energy is mainly
absorbed in the atmosphere by molecular oxygen (O2), ozone
(O3), and water vapor (H2O), as described in Chapter 3.
The absorption of solar radiation by ozone provides the energy that heats the
stratosphere and mesosphere (see Chapter 2). The portion of solar radiation
that is not absorbed in the atmosphere or backscattered to space reaches the
Earth's surface. Figure 15.1 presents the
spectrum of solar radiation outside the Earth's atmosphere and at sea level
for clear sky conditions. Since the troposphere and the surface are coupled
by convective exchanges, this energy almost simultaneously heats the soil,
the vegetation, and the oceans as well as the entire troposphere, except in
cases of temperature inversions near the surface (temperature increasing with
height).

The intensity of radiation emitted by the Sun is not entirely constant as a
function of time. For example, variations in the "solar constant" of
approximately a tenth of a percent are observed and linked to the 11-year
solar cycle. Several attempts have been made to correlate the past evolution
of climate with solar activity, but the subject remains controversial, since
there are no reliable measurements of solar constant changes before the
mid-1970s, and very little change has occurred since then.

15.2.2 Terrestrial Radiation

The energy provided by the Sun and absorbed by the Earth is reradiated as
infrared radiation (see Box 15.2). This energy
is absorbed by clouds as well as atmospheric molecules, the major absorbers
being water vapor and carbon dioxide; these two gases are sufficiently
abundant to trap a large fraction of the energy (mostly in the 12 to 20
µm spectral region) in the lowest layers of the atmosphere. In the 8 to
12 µm region, called the atmospheric window, terrestrial radiation
propagates to space because of the relatively weak absorption in this region
of the spectrum. Therefore, any gas with strong absorption properties in this
spectral region is expected to be relatively efficient in trapping terrestrial
radiation. The spectral locations of the absorption features of the main
greenhouse gases in the atmospheric window are shown in
Figure 15.2. An example of a terrestrial
radiation spectrum measured at the top of the atmosphere by the Nimbus-3 IRIS
instrument is shown in Figure 15.3. The
absorbing bands such as the 9.6 µm band of O3 and the 15
µm band of CO2, as well as the atmospheric window and several
other features (H2O, CH4), are noticeable. These
radiatively active gases, also called greenhouse gases, absorb only a small
fraction of solar energy, but they are very effective in absorbing as well as
emitting longwave radiation. Their net effect is to reduce the amount of
radiative energy emitted to space and to increase the radiative energy
provided to the surface-troposphere system. The fundamental reason for the
existence of the "greenhouse effect" is that the temperature decreases with
altitude in the troposphere. Radiatively active gases as well as clouds
absorb the radiation emitted by the warmer surface, while their emission of
radiation to space occurs at colder atmospheric temperatures. The trapping of
the radiation by radiatively active molecules produces an increase in the
surface temperature of about 33 degrees Celsius (assuming no change in albedo
when atmosphere is removed). Without the "greenhouse effect" the average
temperature at the surface would be only -18°C and life would not be
possible on Earth. At higher altitudes, the radiative emission to space in the
15 µm band of CO2 contributes to a cooling in the stratosphere
and mesosphere.

Surface climates are directly influenced by the radiation balance between
incoming solar radiation and outgoing (reflected solar + infrared) radiation.
A global energy budget of the Earth system can be approximately described as
follows (Fig. 15.4): The solar energy
penetrating into the Earth's system represents about 342 W m-2, of
which about 107 W m-2 (or 31%) is returned to space (24% due to
backscattering by clouds, air molecules, and particles, and 7% due to
reflection at the Earth's surface), 87 W m-2 (or 25%) is absorbed
within the atmosphere, by ozone in the stratosphere, and by clouds and water
in the troposphere. The remaining 148 W m-2 (or 43%) is absorbed at
the Earth's surface. From the terrestrial energy emitted by the surface (390
W m-2 or 114%), only 40 W m-2 (12% of the incoming solar
radiation) escapes directly to space in the atmospheric window. The remaining
310 W m-2 (or 90%) is absorbed within the troposphere by water
vapor, CO2, O3, and the other greenhouse gases, as well
as by clouds and aerosols. Finally, an energy of about 335 W m-2
(or 98%) is emitted back to the surface, while 195 W m-2 (or 57%)
is emitted to space. The excess energy received by the surface is compensated
by nonradiative processes such as evaporation (latent heat flux of 78 W
m-2, or 23%) and turbulence (sensible heat flux of 15 W
m-2, or 4%).

Note the difference between the radiative emission at the Earth's surface
(390 W m-2) and the total infrared emission to space (40 + 195 =
235 W m-2). This energy trapped in the atmosphere (155
W m-2) represents the greenhouse effect. With the exception of
numbers given for the top of the atmosphere, these values are uncertain by
approximately 10 to 20%.

Among the gases present in the atmosphere, the largest contribution to the
greenhouse effect is provided by water vapor, followed by CO2 and
other trace gases such as CH4, N2O, O3, CFCs,
HCFCs, and HFCs.

Clouds also absorb and emit infrared radiation. In addition, they increase
the planetary albedo (defined as the fraction of the incoming radiation that
is reflected). Their net effect on the climate system is complex. It is
believed that high-altitude cirrus clouds contribute to warming, while
low-level stratus clouds contribute to cooling. Overall, the presence of
clouds tends to cool the Earth's system. However, complex feedback effects
could lead to warming or cooling effects, depending on specific cloud changes
in response to changes in climate forcing.

As the concentration of several radiatively active gases are increasing as a
result of human activities, there is great concern about a possible increase
in the greenhouse forcing. When the concentration of a radiatively active
gas increases, initially the longwave radiation to space is reduced. As a
result, the energy budget is out of balance at the top of the atmosphere. At
the same time, if this gas does not affect the absorption of solar radiation,
the net radiative energy available in the lower atmosphere and at the Earth's
surface increases and the energy balance is restored through a warming
of the surface-troposphere system.

15.3 Natural Variations: Past Climates

One of the most fascinating confirmations of a link between atmospheric trace
gas concentrations and climate is provided by the data obtained from ice cores
collected in polar regions. Figure 15.5 shows
the correlations between atmospheric trace gas content and temperature over
the past 240,000 years. Clearly (assuming that the Vostok temperature series
is representative of the global mean temperature), when the atmospheric
abundances of CO2 and CH4 are low, the Earth is in a
relatively cool climate state. To first order, when greenhouse gases are in
relatively low abundance in the atmosphere, there is less infrared trapping of
heat and the Earth surface is cool, perhaps cool enough to initiate
glaciation. However, it should be noted that it is not clear if CO2
was lower before or after the cooling occurred. Phase shifting of the
temperature changes and atmospheric CH4 and CO2
concentrations is evident in some data sets. It is not straightforward to
assign cause and effect relationships based on these time lags due to possible
amplification effects (feedback mechanisms in the Earth system). Changes in
surface temperatures and the areas covered by ice sheet have undoubtedly
affected the exchange rates of greenhouse gases between the surface
(continents, oceans) and the atmosphere. Simultaneously, the atmospheric
abundance of these gases has affected the Earth's climate.

On even longer time scales, the atmospheric concentration of
CO2 may have been as high as 3000 ppmv during the Cretaceous period
(90-65 million years ago). The Earth climate system was, based on geological
evidence, very warm during this period. The area of the polar ice caps was
substantially reduced, and probably absent, and the sea level was much higher
than during present-day conditions. One of the most puzzling aspects of
long-term changes in climate and the abundance of atmospheric greenhouse gases
is what initiates a shift in dominant climate. A postulated cause is provided
by periodic changes in the relative orientation of the Earth to the Sun. This
effect is known as Milankovich forcing. Spectral analyses of geological
records have highlighted typical periods of 20,000, 40,000, and 100,000 years
in climate changes. The observed evolution in the temperature over the past
millenia may have resulted from interactions between such external forcings
(associated with changes in the orbital parameters of the Earth) and internal
dynamics (e.g., glacial feedback response).

It is interesting to note that the past 10,000 years in the history of our
planet, during which civilization developed, were warm and marked by
relatively consistent weather. The previous millenium, a long, cold period
that ended about 11,000 years ago, is known as the "Younger Dryas" with low
CO2 and CH4 concentrations. Recent analyses of ice
cores extracted from the Greenland ice cap suggest that the winter temperature
in Northern Europe has fluctuated by as much as 10°C over periods of time
as short as a decade. These abrupt shifts, which may be related to
re-organizations in the ocean's circulation, could be a manifestation of
chaotic behavior of the climate system (Broecker, 1995). This situation
probably limits our ability to predict future long-term climate evolution.

15.4 Impact of Anthropogenic Trace Gases on Climate

Although water vapor is the most important greenhouse gas, its distribution in
the atmosphere is mainly driven by physical processes, but is only slightly
affected by human activities (e.g., deforestation on a large scale,
which can matter regionally). Human activities are, however, responsible for
significant changes in the abundance of other radiative gases, as seen in
previous chapters.

Table 15.1 provides an estimate
of the greenhouse gas concentrations and trends; the origin and magnitude of
these trends are further discussed in Chapters 5, 8, and 9.

Changes in the distributions of the trace gases are expected to affect the
climate system through spatial and temporal changes in the flux of radiative
energy into and out of the surface-troposphere system. This impact can, in
principle, be quantified by calculating the induced change in the surface
temperature; however, this quantity is largely dependent on complex feedback
processes, which are not fully understood nor easily represented or verified
in existing climate models. It is more straightforward to calculate the
radiative forcing. This quantity is defined as the response in the net
radiative energy flux at the tropopause to changes in the concentration of a
given trace gas. Several factors determine the ability of an atmospheric gas
to affect the radiative forcing: its atmospheric concentration, the strength
and spectral position of its absorption bands, temperature, and pressure.

The radiative forcing associated with changes in the abundance of trace gases
can be estimated by radiative models. The solution of the radiative transfer
equations is complex since the absorption spectra of atmospheric molecules
exhibit structures characterizing their numerous rotation and
vibration-rotation lines (see Box 15.2). The
absorption by these molecules varies considerably over small wavelength
regions, and exact calculations require line-by-line integrations. However,
such calculations require very large amounts of computer time. This direct
approach is usually replaced by an approximate method in which the line or
band characteristics are expressed in terms of global parameters. Radiative
models have been developed to treat either sections of bands (narrow-band
model) or entire bands (wide-band models) (Tiwari, 1978). The radiative codes
included in multidimensional models (2D and 3D) use generally one of these two
approaches.

15.4.1 Direct Radiative Effects

CH4, O3, N2O, CFC-11, CFC-12, and various
CFCs and HCFCs have strong absorption bands in the atmospheric window region.
These trace gases absorb and emit radiation in bands composed of discrete
lines with extended wings. For gases that are present in small quantities,
such as the CFCs and HCFCs, the absorption increases quasilinearly with their
atmospheric concentration. However, for gases with larger concentrations, such
as methane and nitrous oxide, the absorption at the center of the bands is
already saturated for present atmospheric abundances, and increasing
absorption with increasing concentrations occurs mainly in the wings of the
lines. In this case the radiative forcing is approximately proportional to the
square root of the concentration change. For the most abundant radiatively
active species such as CO2, the atmosphere is almost entirely
opaque in the center of the absorption lines, and the radiative effect of
adding CO2 is only noticeable in the wings of the lines. In this
case, the absorption can be approximated by a logarithmic relationship with
the CO2 concentration. A doubling in the atmospheric abundance of
CO2 leads to an increase in the radiative forcing of about 4.6 W
m-2.

Simplified expressions providing the direct radiative forcing as a function of
the greenhouse gases concentrations are given in
Table 15.2 for both clear sky and average
cloudiness conditions. These formulas are only approximate and should be
applied only for small changes in the concentrations.

A useful index also used to evaluate the radiative impact of increasing
greenhouse gases concentrations is provided by the relative radiative forcing.
This index provides a direct comparison between the direct radiative forcing
of a greenhouse gas and that of CO2 (chosen here as the reference
molecule). Values of the direct relative radiative forcing for several
greenhouse gases are given in Table 15.3
for a 1990 reference atmosphere, for both clear sky and cloudy conditions,
calculated for a 10% increase in the concentrations of the more abundant
greenhouse gases. It should be emphasized that since radiative fluxes do not
change linearly with the concentrations of trace gases, the calculated
relative radiative forcing depends on the background concentrations for the
greenhouse gases, as well as the magnitude of the changes applied to these
concentrations. As an example, the relative forcing of methane as compared to
CO2 would be 24.4 for an increase of 1% in the concentrations of
both CO2 and CH4, while it would be 27.6 for a doubling
in the concentrations of both CH4 and CO2.

15.4.2 Indirect Effects: Chemical Feedbacks

Many climate models including general circulation models (GCMs), which are
used to predict climate changes, use CO2 as a proxy for other
greenhouse gases and often estimate climate changes for a doubling in the
equivalent CO2 concentrations. Most of them currently do not
explicitly account for the greenhouse effect of other trace gases. The use of
a CO2 proxy to represent the combined greenhouse forcing of
CO2 and the other radiatively active trace gases is questionable
due to the differences in the spectral and chemical properties of all gases
involved. For example, Figure 15.6 shows the
difference in the heating rate calculated for a doubling of the CO2
concentration (Fig. 15.6a) and an explicit
treatment of the increase in the other trace gases
(Fig. 15.6b). Compared with a doubling of the
CO2 concentration, the inclusion of the radiative effect of the
other trace gases results in a value of the heating rate stronger in the lower
stratosphere and much lower at the surface for high latitudes.

Gases such as CH4, N2O, and the CFCs are not only
radiatively active, but they also produce chemical perturbations in the
atmosphere and hence affect the abundance of other greenhouse gases. The
oxidation of methane, for example, leads to an additional production of water
vapor in the stratosphere and ozone in the troposphere. The breakdown of
N2O and CFCs in the stratosphere leads to the production of active
nitrogen or chlorine radicals that destroy ozone. CO2 is chemically
inactive in the atmosphere, but an increase in the CO2
concentration and in the associated emission to space of the 15 µm
radiation is expected to produce a cooling of the stratosphere and mesosphere.
As the production and loss rates of ozone are strongly temperature dependent
in the middle atmosphere, a CO2 increase is expected to moderate
the ozone destruction caused by chlorine and nitrogen compounds. Furthermore,
a cooling of the winter polar stratosphere resulting from increasing
CO2 concentrations could lead to the formation of additional polar
stratospheric clouds, which are associated with the observed dramatic
destruction of ozone in the Antarctic polar stratosphere.

As a result of all these processes, stratospheric ozone could decrease
globally, and more solar radiation could become available in the troposphere,
leading to a warming of the surface. Less terrestrial radiation would be
absorbed by ozone, leading in this case to a cooling of the surface. The net
effect of stratospheric ozone changes on the climate system depends strongly
on the magnitude and altitude of these changes. Moreover, all these changes
could induce a modification of the circulation in the stratosphere, and thus
affect the transport of other trace gases.

Human activities could also lead to an increase in the concentrations of ozone
in the troposphere with potential impact (absorption of solar and terrestrial
radiation) on the climate system. Such changes probably have a larger impact
on the radiative forcing than those produced by ozone depletion in the
stratosphere.

Increased emissions of trace gases at the Earth's surface could have a
significant impact on the climate system (temperature, precipitation,
frequency of extreme events), but the resulting effects are difficult to
quantify because of strong non-linearities in the coupled chemical and climate
systems. The available estimates of potential climate changes produced by
perturbations in the chemical composition of the atmosphere are provided by
interactive chemical-radiative-dynamical models.

The importance of these chemical feedbacks on the radiative forcing of the
atmosphere is illustrated in Table
15.3. The indirect relative forcing of greenhouse gases has been derived
by using an interactive two-dimensional model that is run to steady state. As
for the calculations of the direct radiative forcing, a 1990 reference
atmosphere is assumed; for each individual gas, a 10% increase in the
background concentration is applied at the surface level.

Another illustration of the importance of chemical processes on the climate
forcing is given in Figure 15.7. This figure
represents the changes from 1900 to 1990 in the radiative forcing (1) when
only considering direct radiative effects and (2) when chemical feedbacks are
taken into account. The changes in the concentration of the greenhouse gases
at the surface from 1900 to 1990 used in these calculations is given in
Table 15.1. When chemical feedbacks are
taken into account, the ozone produced as a result of the CH4
release has a significant radiative effect, which is as strong as the direct
radiative effect of methane.

15.5 Global Warming Potentials (GWPs)

The radiative forcing provides an estimate of the change in the radiative
flux at the tropopause in response to changes in the concentration of
greenhouse gases. In order to take into account the lifetime of the gases in
the atmosphere, and hence the period of time over which the climatic effect
of a perturbation in their concentration is expected to be significant, an
index called the Global Warming Potential (GWP) was defined. This concept was
created in order to enable decision makers to evaluate options to regulate
future emissions of various greenhouse gases without having to perform complex
model calculations.

The GWP of a well-mixed gas is defined (IPCC, 1990) as the time-integrated
change in the radiative forcing due to the instantaneous release of 1 kg of
a trace gas i expressed relative to that from the release of 1 kg of
CO2

GWP =

integralT0
Delta FR, i (t)
dt integralT0 Delta FR, CO2 (t) dt

(15.1)

if DeltaFR represents the
change in the forcing at the tropopause and T is the time over which
the integration is performed (time horizon).
Using a linear approximation,

GWP =

integralT0ai ni (t) dt integralT0aCO2nCO2(t) dt

(15.2)

where ai (expressed in W m-2 kg-1) is
the instantaneous radiative forcing due to the increase in the concentration
of trace gas i and ni is the concentration of the gas
i remaining at time t after the release (IPCC, 1990).
aCO2 and nCO2 are
the corresponding variables applied to CO2, which is considered the
reference gas. If taui is the lifetime
of the molecule i and tau an "effective"
residence time for CO2, the GWP of the gas i can be
approximated by

GWP =

aiintegralT0e-t/
tauidtaCO2integralT0e-t/
tauCO2dt

=

aitauiaCO2tauCO2

1 - e-T/
taui1 - e-T/
tauCO2

(15.3)

As indicated in the above expression, the estimation of the GWP for a trace
gas requires estimates of the radiative forcing for the trace gas i and
for the reference gas CO2 per unit of mass change, the lifetimes of
species i and of CO2, and the definition of the time horizon
T over which the integration is performed. The indirect chemical
effects resulting from the increase in the concentration of species i
also need to be evaluated.

The choice of the time horizon T depends on the type of climate impact
under consideration. As each response has its own characteristic time, there
is no single universally accepted value of the time horizon T that can
be adopted. Table 15.4 illustrates the
integration periods that are appropriate for different climatic responses.

As discussed in Chapter 5, the atmospheric abundance of CO2 is
regulated by the cycling of carbon between several biogeochemical reservoirs
(atmosphere, ocean, biosphere). A single global residence time of
CO2 in the atmosphere cannot be derived. Carbon dioxide added to
the atmosphere decays relatively rapidly over the first 10 years, with a more
gradual decay over the next 100 years and a very slow decline over the 1000
year time scale. Expressions have been deduced from ocean-atmosphere-biosphere
models that provide the decay of a perturbation in atmospheric CO2
as a function of time. The study by Maier-Reimer and Hasselmann (1987), for
example, provides an effective residence time of approximately 120 years for
atmospheric CO2. Typical global lifetimes of other trace gases are
given in Table 15.5.

Table 15.5 presents a recent estimate of
GWPs provided by IPCC (1995). GWPs were calculated by injecting a finite
amount of trace gases to the abundance of the background atmosphere and by
calculating the radiative response over several time horizons (20, 100, 500
years). The model of Siegenthaler and Joos (1992) was used to estimate the
decay response of CO2. For these calculations, the background
atmospheric trace gas concentrations were held fixed (at current levels) and
did not account for a possible future evolution of the atmospheric
composition. Note that the decay time of a methane pulse (12-18 years) is
higher than the global lifetime of this gas, since the concentration of OH
decreases as CH4 is added to the atmosphere, and the resulting GWP
for CH4 is higher than the direct GWP. The estimate of all
indirect effects on the GWP (e.g., changes in ozone, water vapor,
temperature, etc.) is not straightforward, and is generally model dependent.

15.6 Radiative Effects of Aerosols15.6.1 Direct Effects

Aerosols present in the atmosphere (including sulfate particles resulting
from fossil fuel combustion and elemental carbon, EC, released by biomass
burning; see Color Plate 12) absorb and scatter a significant fraction of
incoming solar radiation back to space. The addition of anthropogenic (not
EC) and volcanic aerosols to the atmosphere leads therefore to a reduction in
the net radiation available at the surface and so to a cooling of the Earth's
system. Aerosols also absorb terrestrial radiation and thereby produce a
significant heating in dense aerosol layers.

An interesting modeling study of direct forcing by sulfate aerosols and
comparison with greenhouse gas forcing has been presented by Kiehl and
Briegleb (1993). Best estimates from observations and modeling studies of
sulfate aerosol loadings and distributions were combined with imposed
lognormal aerosol distributions (as discussed in Chapter 4) and derived
optical properties, and the sensitivity of their findings to these aerosol
parameters were examined (Fig. 15.8). They
estimate that variations in size or chemical composition would alter the
estimated forcing (-0.3 W m-2, annually averaged) by ±10%.
The spatial distribution of aerosol properties may have a larger effect. This
is because the greenhouse gas forcing occurs in different regions of the globe
than does the anthropogenic aerosol forcing, which is strongest in the
midlatitudes of the Northern Hemisphere, where most of the sources are
located. In contrast, greenhouse gases, except ozone, generally become well
mixed in the troposphere; their radiative effects are strongest in the region
between -30° and +30° latitude. The combined effects of aerosols
and greenhouse gases thus do not "cancel," but may change global temperature
gradients.

It is known that organic matter can comprise a significant fraction of the
tropospheric aerosol, and thus must also have a role in postulated climate
effects. There is also substantial evidence that some of these species are
hygroscopic and thus should contribute to indirect climate effects as well
(Novakov and Penner, 1993). Soot has been detected in all regions of the
globe, even in "remote" areas. Its strong solar radiation absorption
characteristics suggest that climate forcing due to suspended soot aerosol
will have a sign opposite that of sulfate. The net radiative forcing of a
mixture of sulfate and soot could therefore be substantially smaller than the
forcing calculated for sulfate only.

Widespread dust plumes are often detected in satellite images, and it might
be expected that dust contributes to aerosol radiative forcing. Mineral dust
aerosol may both scatter and absorb solar radiation, depending upon its
composition and the wavelength of light considered. Sokolik et al.
(1993) compared measurements of the complex refractive index for atmospheric
dust aerosols and showed that the large range of values for the imaginary part
of the refractive index leads to significant differences in estimates of
radiative forcing. The uncertainty is magnified when one considers the
effects of the presence of other suspended material (e.g., soot).
Rather than dust inducing a significant effect on climate, the major impacts
may follow in the opposite direction; that is, climate change may
significantly affect dust production and transport. Changes in aridity in
North Africa and shifts in large-scale atmospheric circulation patterns
associated with climate change may alter the magnitude and pattern of Saharan
dust transport to the North Atlantic (Arimoto et al., 1992).

Large volcanic eruptions, such as those of El Chichón (1982) and Mt.
Pinatubo (1991), have substantially enhanced the aerosol load of the
atmosphere for a few years, resulting in a noticeable cooling of the surface.
One year after the eruption of Mt. Pinatubo the radiation forcing was
estimated to have been -4 W m-2 while an anomaly of -0.3 to
-0.4°C in the global temperature was reported (Dutton and Christy, 1992;
IPCC, 1995). Such volcanic perturbations are, however, transitory, with a
typical time constant of 1-2 years.

The radiative forcing produced by the enhanced anthropogenic sulfate aerosol
burden since preindustrial times is estimated (on the global scale) to be
approximately -0.6±0.3 W m-2 (IPCC, 1995). Because of
their relatively limited lifetime (a few days), anthropogenic aerosols are
mostly concentrated in industrialized regions (eastern United States, Europe,
eastern Asia), where their radiative impact is believed to be significant.
Although on the global scale their radiative impact is considerably
smaller than the forcing caused by anthropogenic greenhouse gases
(approximately 2.5 W m-2), in industrialized areas the cooling
caused by aerosols exceeds the warming produced by enhanced CO2 and
other radiatively active gases (Fig. 15.9).

Sulfate aerosols also serve as cloud condensation nuclei (CCN) and hence
affect the formation and the radiative properties of clouds. This indirect
climate impact of anthropogenic sulfur remains poorly quantified, but could
be as large as or even larger than the direct forcing by aerosols of human
origin.

An understanding of the role of aerosol mixtures in the past and for the
present day, and extensions to predict climate change, is hindered by large
uncertainties in many key quantities needed for such estimates.
Quantification of major uncertainties and a proposal for strategies for
minimizing them are presented in the review by Penner et al. (1993)
(Table 15.6).

15.6.2 Indirect Effects

The subset of atmospheric aerosols active as CCN may have an "indirect effect"
on climate by altering the albedo of clouds. Changes in the availability of
CCN may change nucleated droplet number concentrations. As droplet number
concentrations increase for fixed liquid water content, the mean droplet size
decreases and the reflectivity of the cloud increases; the global energy
balance is sensitive to such changes. Studies of potential "indirect" aerosol
effects have focused upon the role of marine stratocumulus clouds, in part
because of their ubiquity (they cover about 25% of the Earth's surface) and
because the potential for perturbations to these clouds is high. Marine
clouds generally have low droplet concentrations (on the order of 100
cm-3), believed to be limited by the availability of CCN. Any
process, then, that alters the relative amounts of CCN in marine regions may
affect the albedo of these clouds. In contrast, clouds formed over
continental regions are believed to have an excess of CCN available, and the
number activated is most likely related to other factors such as the maximum
supersaturation. However, there is observational evidence for a dependence
of continental cloud drop concentration on aerosol loading (e.g.,
Leaitch et al., 1992). Han et al. (1994) derived effective
cloud drop radii from satellite data and reported systematic differences in
drop size between continental and marine water clouds and between marine
clouds in the Northern and in the Southern Hemispheres. Smaller drop radii
were found in those regions most affected by anthropogenic pollution, in
support of the "indirect effect" hypothesis.

The importance of sulfate to aerosol and CCN concentrations led to the
interesting hypothesis of Charlson et al. (1987) of a climate feedback
loop involving marine phytoplankton, DMS emissions, CCN concentrations, and
cloud radiative forcing. This work spurred much of the subsequent research
and debate regarding indirect climate effects of aerosols. Some estimates
suggest that a 30% change in CCN available to marine stratus will lead to a
globally averaged forcing of 1 W m-2. However, changes in CCN
populations may have other effects that also influence climate. For example,
enhanced droplet concentrations may reduce the likelihood of precipitation
from clouds, altering cloud cover and cloud lifetime (Radke et al.,
1989). The response of cloud liquid water content to changes in CCN and
climate is not well understood. Changes in precipitation would also change
the atmospheric concentration of the most important greenhouse gas: water
vapor.

Climate effects of aerosols are also postulated for polar regions. The
climate of polar regions is of great interest in studies of global warming.
As temperatures increase the extent of snow and ice is reduced, decreasing the
surface albedo and further increasing the amount of sunlight that is absorbed
by the Earth-atmosphere system. Conversely, a temperature decrease will
increase the surface albedo and thus reinforce the cooling (e.g.,
Curry et al., 1993). This feedback mechanism results in the Arctic
having an impact on the global climate as well as the local climate, since the
ice-albedo feedback mechanism can result in substantial modification of the
net energy retained by the Earth-atmosphere system.

Several types of polar aerosol effects are postulated. Soot aerosols
deposited to snow and ice surfaces may alter their albedo in a "direct"
effect. An "indirect effect" for polar ice-phase clouds (present even in
the lower troposphere during the coldest months of the year) may also occur
via the following mechanism. Polluted air has been typically shown to be
deficient in ice-forming nuclei (IN). This relationship is believed to arise
from an increased sulfate mass loading in polluted air; sulfate particles,
which are poor IN, coagulate with potential IN and effectively deactivate
them (Borys, 1989). If this hypothesis is correct, ice nucleation in the
Arctic may be relatively enhanced during winter, if there is a decrease in the
oxidation of SO2 in the relative absence of sunlight and liquid water,
which could result in a decreased amount of sulfate particles. Conversely,
ice nucleation during "Arctic haze" events in the Spring would be suppressed.
Thus, anthropogenic aerosol has the potential to impact the amount of
condensed water and the total water budget in the Arctic by modifying the ice
nucleation and the phase of condensed water.

15.7 Response of the Climate System to Radiative Forcing

The simplest model to estimate the response of the Earth's climate to
radiative perturbations is based on the global balance between the incoming
solar energy (FS) that is absorbed by the Earth system and
the outgoing terrestrial radiative energy (FT) that is
emitted to space. For the system to be at equilibrium

FS = FT
(15.4)

If the system is perturbed by some radiative forcing
DeltaFR (e.g., by an increase in the
atmospheric abundance of carbon dioxide), the equilibrium between incoming and
outgoing energy will be restored by a change (Delta
) in the initial fluxes (FS and FT),
such that

Delta (FT - FS) =
DeltaFR
(15.5)

Assuming that the balance is re-established by a change in surface temperature
(DeltaTs), the climate
sensitivity factorlambdac, defined
by

DeltaTs =
lambdacDeltaFR
(15.6)

is simply given by

lambdac = (
DeltaFT / DeltaTs - DeltaFS
/ DeltaTs)-1
(15.7)

To make a first-order estimate of this factor, we assume that the globally
averaged solar energy absorbed by the Earth system is given by

FS = F0 / 4 (1 -
alpha)
(15.8)

where alpha is the Earth's albedo (typically 0.3)
and F0 = 1370 W m-2 is the solar constant (the
averaged solar energy intercepted by a sphere is F0 /4).
Similarly, we assume that the terrestrial energy radiated to space is
expressed by the Stefan-Boltzmann law

FT = epsilon sigmaTs4
(15.9)

where sigma is the Stefan-Boltzmann constant,
epsilon the emissivity of the atmosphere, and
Ts the surface temperature. Assuming that
alpha and epsilon are
constant, and neglecting all potential feedbacks in the climate system, one
derives from Eqs. (15.7-15.9) that

lambdac = (4
epsilon sigmaTs3)
-1 = Ts / 4FT
(15.10)

Under these assumptions, the value of the climate sensitivity factor is
approximately equal to 0.3 K (W m-2)-1 (Kiehl, 1992).
Thus, for a perturbation associated with a doubling in the CO2
abundance (DeltaFR = 4.6
W m-2), the increase DeltaTs in the mean Earth's temperature is 1.4 K, that is,
significantly less than predicted by climate models or derived from satellite
observations [lambdac
~/= 0.6 K (W m-2)-1].

The reason for this discrepancy is that important climate feedbacks have been
ignored in this simple calculation. Examples of such feedbacks include those
associated with the hydrological cycle. As a result of enhanced radiative
forcing, the warmer atmosphere becomes more humid (Clausius-Clapeyron
relation, e.g., see Hartmann, 1994), which produces an additional
greenhouse forcing and hence a larger warming of the planet. Another positive
feedback is produced by changes in the surface albedo
[alpha in Eq. (15.8)] when the surface area
covered by ice and snow varies in response to climate change. Feedbacks
caused by changes in cloudiness in response to global warming/cooling are
difficult to estimate since clouds reflect solar radiation back to space
(cooling effect) and, at the same time, reduce the emission to space of
terrestrial radiation (warming effect). Finally, the assessment of potential
feedbacks involving the biosphere remains an important research topic.

Modern climate models account for many of the relevant feedback mechanisms.
When used in a predictive mode, these models attempt to simulate the transient
response of the climate system to changes in the radiative forcing. This
transient response is determined by the thermal inertia of the system, that
is, the effective heat capacity of the atmosphere, land, and ocean, as well as
the radiative damping of the system (Schneider, 1992). The response of
climate to a gradual increase in the atmospheric abundance of radiatively
active gases can therefore be modeled accurately only with coupled
ocean-atmosphere models. A representation of a complex climate model is shown
in Figure 15.10. Such models show that the
existence of feedbacks associated with water vapor, clouds, and ice tends to
increase the value of the
lambdac factor to a value
ranging between 0.4 and 1.25 K (W m-2)-1 depending on
the formulation in the model of the feedback processes. Thus, with these
values of the feedback factor, the warming (at equilibrium) caused by a
doubling in the atmospheric concentration of CO2 should range from
1.8 to 5.8 K. The change in temperature for a CO2 doubling,
however, is not uniform in space and, as shown by
Figure 15.11, is expected to be largest at
high latitudes in winter. Improvements in climate predictions require,
therefore, that physical, chemical, and biological processes be better
understood and more accurately represented in the models.

Further Reading

Brasseur, G., ed. (1997) The Stratosphere and Its Role is the Climate
System, NATO ASI Series I-54, Springer-Verlag, Berlin.